American Mineralogist, Volume 100, pages 824–836, 2015
The origin and implications of clay minerals from Yellowknife Bay, Gale crater, Mars†k
Thomas F. Bristow1,*, David L. Bish2, David T. Vaniman3, Richard V. Morris4, David F. Blake1,
John P. Grotzinger5, Elizabeth B. Rampe4, Joy A. Crisp6, Cherie N. Achilles2, Doug W. Ming4,
Bethany L. Ehlmann5,6, Penelope L. King7,8, John C. Bridges9, Jennifer L. Eigenbrode10,
Dawn Y. Sumner11, Steve J. Chipera12, John Michael Moorokian6, Allan H. Treiman13,
Shaunna M. Morrison14, Robert T. Downs14, Jack D. Farmer15, David Des Marais1,
Philippe Sarrazin16, Melissa M. Floyd10, Michael A. Mischna6 and Amy C. McAdam10
Exobiology Branch, NASA Ames Research Center, Moffett Field, California 94035, U.S.A.
Department of Geological Sciences, Indiana University, 1001 East Tenth Street, Bloomington, Indiana, 47405, U.S.A.
3
Planetary Science Institute, 1700 E. Fort Lowell, Tucson, Arizona 85719-2395, U.S.A.
4
ARES Division, NASA Johnson Space Center, Houston, Texas 77058, U.S.A.
5
Division of Geological and Planetary Sciences, California Institute of Technology, 1200 E. California Boulevard, Pasadena, California 91125, U.S.A.
6
Jet Propulsion Laboratory, California Institute of Technology, 4800 Oak Grove Drive, Pasadena, California 91109, U.S.A.
7
Research School of Earth Sciences, Australian National University, Canberra ACT 0200, Australia
8
Department of Physics, University of Guelph, Guelph, Ontario N1G 2W1, Canada
9
Space Research Center, University of Leicester, Leicester LE1 7RH, U.K.
10
NASA Goddard Space Flight Center, Greenbelt, Maryland 20771, U.S.A.
11
Department of Earth and Planetary Sciences, University of California, Davis, California 95616, U.S.A.
12
Chesapeake Energy Corporation, 6100 N. Western Avenue, Oklahoma City, Oklahoma 73118, U.S.A.
13
Lunar and Planetary Institute, 3600 Bay Area Boulevard, Houston, Texas 77058, U.S.A.
14
Department of Geology, University of Arizona, Tucson, Arizona 85721, U.S.A.
15
Department of Geological Sciences, Arizona State University, Tempe, Arizona 85281, U.S.A.
16
SETI Institute, Mountain View, California 94043, U.S.A.
1
2
Abstract
The Mars Science Laboratory (MSL) rover Curiosity has documented a section of fluvio-lacustrine strata at
Yellowknife Bay (YKB), an embayment on the floor of Gale crater, approximately 500 m east of the Bradbury
landing site. X‑ray diffraction (XRD) data and evolved gas analysis (EGA) data from the CheMin and SAM
instruments show that two powdered mudstone samples (named John Klein and Cumberland) drilled from the
Sheepbed member of this succession contain up to ~20 wt% clay minerals. A trioctahedral smectite, likely a
ferrian saponite, is the only clay mineral phase detected in these samples. Smectites of the two samples exhibit
different 001 spacing under the low partial pressures of H2O inside the CheMin instrument (relative humidity
<1%). Smectite interlayers in John Klein collapsed sometime between clay mineral formation and the time of
analysis to a basal spacing of 10 Å, but largely remain open in the Cumberland sample with a basal spacing of
~13.2 Å. Partial intercalation of Cumberland smectites by metal-hydroxyl groups, a common process in certain
pedogenic and lacustrine settings on Earth, is our favored explanation for these differences.
The relatively low abundances of olivine and enriched levels of magnetite in the Sheepbed mudstone,
when compared with regional basalt compositions derived from orbital data, suggest that clay minerals formed
with magnetite in situ via aqueous alteration of olivine. Mass-balance calculations are permissive of such a
reaction. Moreover, the Sheepbed mudstone mineral assemblage is consistent with minimal inputs of detrital
clay minerals from the crater walls and rim. Early diagenetic fabrics suggest clay mineral formation prior to
lithification. Thermodynamic modeling indicates that the production of authigenic magnetite and saponite at
surficial temperatures requires a moderate supply of oxidants, allowing circum-neutral pH. The kinetics of
olivine alteration suggest the presence of fluids for thousands to hundreds of thousands of years. Mineralogical
evidence of the persistence of benign aqueous conditions at YKB for extended periods indicates a potentially
habitable environment where life could establish itself. Mediated oxidation of Fe2+ in olivine to Fe3+ in magnetite, and perhaps in smectites provided a potential energy source for organisms.
Keywords: Mars, Yellowknife Bay, clay minerals, CheMin, XRD, habitability
Introduction
Over the last decade, orbital mineralogical surveys of the
surface of Mars have revealed the widespread presence of clay
* E-mail: thomas.f.bristow@nasa.gov
†kOpen access: Article available to all readers online.
0003-004X/15/0004–824$05.00/DOI: http://dx.doi.org/10.2138/am-2015-5077CCBYNCND
824
minerals, largely restricted to Noachian age terrains (Poulet et
al. 2005; Bibring et al. 2006; Murchie et al. 2009; Ehlmann et al.
2011). As common aqueous alteration products of the basaltic
igneous rocks that make up much of the martian crust, these
clay minerals may indicate a period in Mars history >3.7 Ga
when conditions were wetter and perhaps warmer than today.
BRISTOW ET AL.: CLAY MINERALS AT YELLOWKNIFE BAY
However, questions remain as to whether clay minerals detected
from space reflect conditions at or near the surface of Mars
or are remnants of subsurface hydrologic systems that were
more active early in Mars’ history (Ehlmann et al. 2011). Thus
establishing the geological context and temporal distribution
of clay-bearing deposits has become central to characterizing
potential habitable environments for life and the abundance
and mobility of water on ancient Mars.
One of the primary science goals of the MSL rover Curiosity
is to study a succession of clay mineral- and sulfate-bearing
strata near the base of a 5 km thick sequence of layered rocks
exposed in the flanks of Aeolis Mons, informally called Mt.
Sharp, a mound that rises from the center of ~150 km diameter
Gale crater (Anderson and Bell 2010; Milliken et al. 2010).
The types of clay minerals and the temporal succession of
clay-bearing strata overlain by sulfate-bearing lithologies bear
a striking resemblance to posited global mineralogical trends
documented from space, which may reflect changes in climate
and the availability of water on early Mars (Grotzinger and
Milliken 2012).
After an initial commissioning phase, Curiosity drove
eastward from Bradbury Landing where it encountered clay
mineral-bearing mudstones at Yellowknife Bay (YKB) ~500
m east of the rover’s landing site (Vaniman et al. 2014). The
mudstones are part of a section of the fluvio-lacustrine YKB
formation, derived from erosion of rocks along the Gale crater
rim, as confirmed by a bulk K-Ar age of 4.21 ± 0.35 billion
years (Farley et al. 2014). However, the stratigraphic relationships between the YKB formation and either the Peace Vallis
fan or the strata of lower Mt. Sharp are uncertain. In either
case, strata of the YKB formation, including the clay mineralbearing Sheepbed member, post-date the Noachian-Hesperian
boundary and are younger than the majority of clay minerals
documented on Mars from orbit (Ehlmann et al. 2011; Vaniman
et al. 2014; Grotzinger et al. 2014). The clay minerals at YKB
are interpreted to have formed in situ and indicate paleoenvironmental conditions within the fluvio-lacustrine system were
potentially habitable for life (Vaniman et al. 2014; Grotzinger
et al. 2014). The age of the deposits implies habitable conditions on Mars persisted, at least locally, for longer than previously thought. This paper provides a focused examination of
the origins of the clay minerals and associated mineralogical
processes, considering their sedimentological/mineralogic
context, regional geology, and the geochemical compositions
of the host sediments.
Sedimentary context
A ~4.75 m thick succession of flat lying, decimeter-scale
sedimentary strata, collectively known as the YKB formation,
is exposed in the western margin of YKB—a local depression
~18 m below the elevation of the Bradbury landing site. The
base of the YKB formation is not exposed and total thickness
is not known. The YKB formation largely consists of basaltic
mudstones and sandstones and is informally subdivided into
the Sheepbed, Gillespie Lake, and Glenelg members, with the
Sheepbed member mudstone the lowest unit exposed (Grotzinger et al. 2014). Members are defined based on differences in
lithology, bedding geometry, sedimentary structures, and lateral
825
extent. Facies analysis supports a fluvio-lacustrine origin for
these strata (Grotzinger et al. 2014).
The Sheepbed member is of particular scientific interest
because evidence points to deposition in a standing body of
water (Grotzinger et al. 2014). Regional mapping indicates a
minimum deposit extent of 4 km2, with a thickness of at least
1.5 m at YKB. The Sheepbed mudstone consists of homogenous, fine-grained sediments (nearly all discernable grains
<50 μm, with the clay mineral content presumably <2 μm) with
a few very thin intercalated beds that resist erosion. Curiosity
carried out a detailed geochemical and textural examination
of a traverse through the Sheepbed. The mudstone contains
early diagenetic features including nodules, hollow nodules
and raised ridges (Grotzinger et al. 2014; Stack et al. 2014;
Siebach et al. 2014). The ridges tend to be elevated in magnesium relative to the surrounding sediments (Léveillé et al.
2014). The bulk chemistry of rocks throughout the member is
similar to average martian crust (McLennan et al. 2014). Together, textures, bulk chemistry, and the sedimentary context of
the Sheepbed are most amenable to interpretation as proximal
lacustrine deposits (Grotzinger et al. 2014).
The Sheepbed member (and other parts of YKB formation)
experienced a period of aqueous alteration post-dating lithification. A network of brittle fractures, up to 8 mm in width and
lacking preferred orientation, cut early diagenetic features and
are filled with light-toned hydrated and dehydrated sulfates
including bassanite and anhydrite (Grotzinger et al. 2014;
Vaniman et al. 2014). These Ca-sulfates fill some of the hollow
nodules that are intersected by the late-diagenetic fractures.
Curiosity drilled two ~6 cm deep holes in the Sheepbed
member at sites named John Klein and Cumberland. The
sites are 2.5 m apart and are within ~30 cm of each other
stratigraphically. John Klein is a representative lithology of
most of the Sheepbed member, whereas Cumberland contains
abundant early diagenetic concretions, especially hollow nodules (Grotzinger et al. 2014; Stack et al. 2014). Drilling of the
Sheepbed member provided images of gray mudstone tailings,
in striking contrast with the dusty, red surface. Drill powders
were delivered to CheMin and SAM, instruments onboard Curiosity that perform X‑ray diffraction and evolved gas analyses,
respectively (Blake et al. 2012; Mahaffy et al. 2012), with the
aim of using mineralogical data and evolved volatile detections
to constrain aqueous conditions during deposition (Vaniman et
al. 2014; Ming et al. 2014). These data also bear on the nature
and influence of post-depositional alteration with implications
for the geological history of Gale and regional stratigraphic
relationships of the sediments at YKB.
Methods: Mineralogical analysis
Curiosity is equipped with a suite of instruments used for direct and indirect
mineralogical analysis (Blake et al. 2012; Campbell et al. 2012; Mahaffy et al.
2012; Wiens et al. 2012). Details of the instruments, data collection, and data
reduction procedures are discussed in online supplemental materials 1, along
with Appendix Figure 11.
Deposit item AM-15-45077, Table 1, Appendix Figure 1, and Methods section.
Deposit items are stored on the MSA web site and available via the American
Mineralogist Table of Contents. Find the article in the table of contents at GSW
(ammin.geoscienceworld.org) or MSA (www.minsocam.org), and then click on
the deposit link.
1
826
BRISTOW ET AL.: CLAY MINERALS AT YELLOWKNIFE BAY
Results
John Klein
Bulk mineralogy
25
20
25
6.44 Å Plagioclase
6.05 Å Bassanite
Intensity
13.2 Å 2:1 clay mineral
Cumberland
5
10
15
2θ
2θ
30
2.53 Å Magnetite
35
2.10 Å Magnetite
3.01 Å Pyroxene
3.22 Å Plagioclase
30
40
35
45
40
50
2.10 Å Magnetite
20
2.53 Å Magnetite
15
3.01 Å Pyroxene
10
4.05 Å Plagioclase
5
6.44 Å Plagioclase
6.05 Å Bassanite
Intensity
10 Å 2:1 clay mineral
John Klein
3.22 Å Plagioclase
4.05 Å Plagioclase
45
50
Figure 1. 1D XRD patterns from John Klein and Cumberland
converted from 2D patterns, with the d-spacing and phase assignments
labeled for major peaks. Diffractogram data were previously presented
in Vaniman et al. (2014).
13.2Å
10.1Å
7.4Å
Intensity
The bulk mineralogy of John Klein and Cumberland has
been discussed by Vaniman et al. (2014) and is summarized
here because it is relevant to discussions of the origins of clay
minerals in the Sheepbed member. CheMin XRD patterns of
both samples show similar crystalline, poorly crystalline phyllosilicate, and X‑ray amorphous components (Fig. 1). The crystalline component largely consists of minerals expected from a
basaltic source (plagioclase feldspar and pyroxenes). Accessory
minerals include metal oxides and Ca-sulfates. Ca-sulfates may
reside entirely in a series of fractures and veins observed in both
drill holes. Fe-forsterite content in John Klein and Cumberland
is 5.7 and 1.9 wt%, respectively, normalized to the crystalline
component excluding clay minerals and the X‑ray amorphous
component. The analyzed Fe-forsterite content of John Klein
may be abnormally high as a result of co-mingling of the John
Klein sample with as much as 5% contamination by residue of
a previous, olivine-rich sample, Rocknest, in the CHIMRA system. In any case, the quantity of Fe-forsterite in both samples is
notably lower than estimates of olivine content of basalts in the
Gale Region based on TES global mapping of 11 wt% (Rogers
and Christensen 2007) and the 22 wt% olivine content of the
nearby Rocknest aeolian sand shadow scooped and analyzed
Cumberland
3
6
9
2θ
12
15
Figure 2. Comparison of John Klein and Cumberland XRD patterns
in the clay mineral 001 region. Labeled basal spacings have been
corrected to for Lorentz polarization.
by Curiosity (Blake et al. 2013; Bish et al. 2013). Furthermore,
the quantities of magnetite in the Sheepbed mudstone are high
(8.7 and 9.5 wt% of the crystalline component at John Klein and
Cumberland, respectively) given that sediments were deposited
from non-turbulent suspensions in a lake as opposed to a vigorous
fluvial setting that could have promoted physical density-driven
sorting (Grotzinger et al. 2014; Vaniman et al. 2014; McLennan
et al. 2014).
Broad 00l diffraction peaks confirm the presence of clay
minerals in John Klein and Cumberland. In the John Klein XRD
pattern, a 001 reflection was observed between 8–11° 2θ CoKα
(9.4 to 12 Å) with a Lorentz polarization corrected peak at ~10.1
Å, indicating the presence of 2:1 group clay minerals (Vaniman et
al. 2014; Fig. 2). The 02l band of clay minerals lies between 22
and 23° 2θ and clearly contributes to the John Klein diffraction
pattern, albeit with overlapping peaks from pigeonite and plagioclase (Vaniman et al. 2014; Fig. 3). A complex 001 peak in the
Cumberland XRD pattern was observed between 6–11° 2θ (9.4
to 17 Å) with a Lorentz polarization corrected primary maxima
at 13.2 Å (Vaniman et al. 2014). A weaker peak is centered at
10° 2θ CoKα (~10 Å) (Fig. 2). Like John Klein, Cumberland 00l
peaks indicate the presence of 2:1 group clay minerals, albeit
with a larger basal spacing (Vaniman et al. 2014). The 02l band
of the Cumberland pattern appears to be in a similar position as
John Klein (~22.5° 2θ, 4.58 Å) (Fig. 3).
The absence of 7 Å peaks in John Klein and Cumberland
excludes the possibility of more than ~5 wt% contributions from
chlorite, kaolin, and serpentine group minerals (detection limit
estimated by Bish et al. 2013). A small peak at 7.4 Å observed
in Cumberland is attributed to akaganeite (Vaniman et al. 2014).
Talc and a disordered talc-like phase (kerolite) are members of
the 2:1 clay mineral group, but they are unlikely to be present
in YKB samples because the 00l peaks do not correspond, respectively, to the 9.38 and 9.65 Å spacing characteristic of these
phases. Furthermore, the breadth of the ~10 Å 001 peak in John
Klein is inconsistent with the presence of well-ordered mica.
The clay mineral content of John Klein and Cumberland
is estimated at 22 and 18 wt%, respectively (±50% relative),
based on whole-pattern fitting of the sample (Vaniman et al.
2014). Independent estimation of clay mineral content based
on SAM EGA measurements of H2O evolved during heating to
BRISTOW ET AL.: CLAY MINERALS AT YELLOWKNIFE BAY
Saponite
b =9.200±0.005Å
Model
Intensity
20
Table 11). No changes were observed in positions of the basal
reflections or other peaks of clay minerals.
a
Plag.
Pigeonite
21
22
2θ
23
24
Saponite
b = 9.204±0.004Å
Saponite
25
b
Intensity
Model
Plag.
Pigeonite
21
22
23
2θ
24
827
Saponite
John Klein
With these results in mind and considering that expandable
clay minerals can collapse to ~10 Å under the low humidity
conditions inside CheMin (Suquet and Pezerat 1987; Sato et al.
1992; Bish et al. 2003; Rampe et al. in preparation), the 00l peak
observed at John Klein is permissive of the presence of discrete
illite, discrete smectite, discrete vermiculite, mixed-layer illitesmectite (I/S), or mixed-layer talc-smectite (T/S).
Chemical constraints preclude illite as the major component
species in John Klein. Mass-balance calculations using bulkrock chemical data from APXS and the chemistry of crystalline
components derived from unit-cell parameters (Vaniman et al.
2014; Morrison, in preparation) indicate a combined K2O content
of the clay mineral and amorphous components in John Klein of
0.69 wt%. K2O content can be used to provide an upper bound
on the discrete illite content of John Klein by assuming the all
the clay mineral and amorphous component K is in discrete illite.
Using a typical structural formula for illite, with 1.5 atoms of K
per formula unit, discrete illite could potentially make up to ~25
wt% of the clay mineral fraction. However, given that much of the
25
Figure 3. 02l region of John Klein (a) and Cumberland (b)
XRD patterns showing data collected (black dots), and the modeled
pattern (solid black line) calculated using Rietveld refinement showing
contributions from saponite, pigeonite, and plagioclase (see description
in text). Refined b unit-cell parameter of saponite is also shown.
temperatures that induce dehydroxylation indicate clay mineral
contents of ~16–17 ± 11–12 wt% (McAdam et al. 2014, and in
preparation).
Correct interpretation of basal reflections of clay minerals
requires consideration of the differences in relative humidity
(RH) on the floor of Gale crater and inside CheMin. Outside the
rover RH ranges from close to 0% at midday to more than 70%
at late night/early morning driven by temperature changes of up
to 60 °C (Fig. 4). Theoretical work on the hydration states of
montmorillonites predicts diurnal changes in interlayer hydration
state under such conditions (Bish et al. 2003). In contrast, RH
remains <1% inside CheMin because of warmer temperatures
are maintained (ranging from 5–25 °C). Laboratory tests of the
time taken for various expandable clay minerals to equilibrate
in CheMin following changes in RH confirm they take place in
a matter of hours to days (Rampe et al. in preparation). This is
longer than similar changes in d-spacings of expandable clay
minerals that occur within minutes when fully exposed to dry
gas in conventional XRD instruments (e.g., Wilson et al. 2004).
CheMin analyzed John Klein and Cumberland at intervals of 13
and 3 sols after drilling, respectively, so these samples should
have attained equilibrium with the gas inside the instrument. To
confirm this and to look for any other mineralogical changes that
may have occurred inside CheMin, John Klein and Cumberland
were re-analyzed at various times after initial analysis (see
Figure 4. Relative humidity (black) and temperature data (red) from
outside the rover (Planetary Data System Data Set ID for the REMS
Instrument: MSL-M-REMS-5-MODRDR-V1.0). Data are averaged over
5 sol periods from sol 196–201 (a) and sol 485–490 (b) showing seasonal
variability. RH inside CheMin, where temperatures are maintained
between 5 to 20 °C remain <1%.
BRISTOW ET AL.: CLAY MINERALS AT YELLOWKNIFE BAY
K could be partitioned into the amorphous material and K may
also reside in exchangeable smectite interlayer sites, this upper
estimate of illite content of John Klein is conservatively high.
The partially ordered stacking of illite gives rise to additional
hkl peaks not seen in collapsed turbostratic smectite >23° 2θ (Co).
It is not possible to evaluate the presence of these directly in the
John Klein (and Cumberland) patterns because of the complexity induced by multiple crystalline phases and the amorphous
component(s) that diffract/scatter in this region. However, wholepattern fitting models of the John Klein XRD pattern including
illite standards (Chipera and Bish 2002) did not improve fits,
permissive evidence for the absence of discrete illite.
The 02l band maximum lies at ~22.5° 2θ (4.58 Å). This is
most consistent with a trioctahedral clay mineral (Moore and
Reynolds 1997), although it is noted that some synthetic Fe
dioctahedral smectites have 02l bands positions close to the
observed position in John Klein (Brigatti 1983; Heuser et al.
2013; Vaniman et al. 2014). Trioctahedral illites are uncommon
on Earth (Galan 2006), and the few examples documented are
transient weathering products of biotite in soils, often associated
with vermiculite (Fordham 1990). Biotite was not detected in the
Sheepbed mudstone and we consider the possibility of a source
of biotite in the catchment remote given the basaltic nature
of the crystalline mineral component observed in John Klein
and Cumberland. Intermediate Mg-rich di-trioctahedral illites
have been reported from ancient terrestrial saline lake deposits
(Hover and Ashley 2003; Deocampo et al. 2009), but based on
deconvolution of the 060 region of XRD patterns of terrestrial
materials their 02l band positions would not match the Sheepbed
clay minerals analyzed by CheMin.
Vermiculite typically forms from chlorite via chemical leaching. Chlorite was not detected by CheMin, and the bulk chemistry
of the Sheepbed mudstone indicates little chemical alteration
of sediments (McLennan et al. 2014), thus arguing against the
presence of vermiculite in John Klein or Cumberland.
Mixed-layer I/S is typically dioctahedral, whereas mixedlayer trioctahedral smectites are usually intergraded with chlorite
(Weaver 1989; Moore and Reynolds 1997). Thus, only small
amounts of I/S are permitted by the chemistry and XRD pattern
data, considered collectively. Another candidate mixed-layer
clay mineral is talc/nontronite, reported from hydrothermally
altered basalts in terrestrial seafloor settings (Cuadros et al.
2013). However, the mixed dioctahedral/trioctahedral nature
of this phase gives rise to 02l bands with two peaks and two
discreet dehydroxylation temperatures (Cuadros et al. 2013) not
supported by current interpretations of XRD and SAM EGA data
(Ming et al. 2014; McAdam et al. 2014, and in preparation).
Our favored interpretation is that the majority of the clay
minerals in the John Klein sample are collapsed trioctahedral
smectite with basal spacing of ~10 Å. Possible candidates
include saponite (with some tetrahedral Al substitution for Si)
and stevensite (which lacks tetrahedral Al substitution for Si).
Unfortunately, precise information on the chemistry of the clay
mineral is not available to make this distinction. However, a
comparison of 02l band positions of clay mineral standards run
on the CheMin 4 instrument, a reasonable proxy for the flight
instrument that is similar in terms of operation and geometry
(Sarrazin et al. 2005), indicates that ferrian saponite, informally
known as Griffith Smectite (Treiman et al. 2015), is the closest
of available materials in terms of 02l peak position. The peak
release of water at ~750 °C is consistent with dehydroxylation
of octahedral sheets of ferrian saponite (Fig. 5; Ming et al. 2014;
McAdam et al. 2014, and in preparation). Additionally, ferrian
saponite fits well with our proposed formation model (see below);
reactive Al from the amorphous component of the Sheepbed
mudstone was likely present during genesis, thus saponite is
predicted (e.g., Bristow and Milliken 2011).
Cumberland
The low-angle region of the Cumberland XRD pattern is
different from John Klein, in that it consists of a broad, complex
peak between 9.4 to 17 Å with the highest intensity centered at
13.2 Å and a smaller rise around 10 Å. The 02l band position
is virtually identical to that in John Klein. However, because
determining the position of the 02l band can be complicated
by interference of pigeonite and plagioclase hkl peaks (Fig. 2),
its position was modeled using BGMN, a Rietveld refinement
program that can generate XRD patterns of partially disordered
clay minerals and consider contributions from other phases
(Bergmann et al. 1998; Ufer et al. 2004). The position of the 02l
band is related to the b unit-cell parameter of the clay mineral
and Rietveld analysis shows the b parameter of clays from John
Klein and Cumberland to be the same within error.
John Klein
Cumberland
b = 9.200 ±0.005 Å
b = 9.204 ±0.004 Å
The chemistry of individual crystalline mineral phases of
John Klein, Cumberland, and the Rocknest sand shadow has
10x104
John Klein
8
Cumberland
Fe-saponite
H2 O counts/s
828
6
4
2
0
0
200
400
Temperature (oC)
600
800
Figure 5. Evolved gas analysis of H2O from a continuous ramp SAM
pyrolysis of drill material from John Klein and Cumberland, compared
with data from a ferrian saponite. The ferrian saponite tested is sample
AMNH 89172 reported in Treiman et al. (2015). EGA data from the
ferrian saponite were acquired under SAM-like conditions in the highfidelity SAM Testbed laboratory instrument suite (McAdam et al. 2014,
and in preparation). Because data from the main mass-to-charge ratio of
H2O (m/z 18, H216O) were saturated, counts from H2O isotopologs and
fragments were plotted to illustrate H2O signal vs. temperature [m/z 20
(H218O) for John Klein and Cumberland and m/z 17 ×0.1 (OH fragment)
for Fe-saponite].
BRISTOW ET AL.: CLAY MINERALS AT YELLOWKNIFE BAY
been constrained based on unit-cell parameters determined by
Rietveld refinement (Bish et al. 2013; Vaniman et al. 2014;
Morrison et al. in preparation). This analysis is made possible
by large databases that allow empirical relationships of mineral
composition and unit-cell parameters to be derived (Morrison et
al. in preparation). Similar attempts have been made to determine
the influence of clay chemistry on the b unit-cell parameter of
trioctahedral smectites based on statistical analysis (Brown and
Brindley 1980). They are related as follows:
b (trioctahedral) = 9.18 + 0.06AlIV – 0.12 AlVI – 0.06Fe3+ + 0.06Fe2+
where AlIV, AlVI, Fe3+, and Fe2+ are the proportions of Al and
Fe atoms per O10(OH)2 in structural formulas, in tetrahedral,
octahedral, ferric, and ferrous oxidation states, respectively. At
the present time, calculated b-parameters for John Klein and
Cumberland cannot be used to constrain detailed clay crystal
chemistry without making major assumptions about the origin
and magnitude of layer charge. However, available XRD data
indicate that the clay minerals in John Klein and Cumberland
are crystallographically similar, albeit exhibiting different
swelling behavior. This similarity is expected given the fact that
Cumberland and John Klein come from the same geological unit
with similar mineralogy and are only 2.5 m apart with ~30 cm
of stratigraphic separation.
Discussion
Swelling behavior
Possible explanations for the difference in 001 position and
swelling behavior between the two Sheepbed samples include
differences in the dominant interlayer cation present in the two
samples and partial intercalation by metal-hydroxyl groups (i.e.,
incipient chloritization) in the Cumberland clay mineral (Vaniman et al. 2014), making it resistant to collapse within the <1%
RH conditions of CheMin sample cells.
Smectites show variable hydration/dehydration properties
depending on the type of interlayer cation as a result of differences in cation hydration energy. The two Sheepbed drill
targets were deliberately chosen because they exhibit different
diagenetic fabrics. John Klein is intensely fractured and filled
with Ca-sulfate-bearing veins, whereas Cumberland is enriched
in early diagenetic veins and nodules (Grotzinger et al. 2014).
Thus, the clay minerals in these two samples may have been
exposed to different fluids (or at least to different degrees) and
may have different interlayer cations.
Of the possible cations found in expandable clays, the divalent
cation Mg2+ is most effective at retaining interlayer H2O. It is
possible that the dominant interlayer cation in Cumberland is
Mg because early diagenetic raised ridges nearby contain Mgsilicates that appear to have precipitated from fluids (Grotzinger
et al. 2014; McLennan et al. 2014; Léveillé et al. 2014). However,
experiments predict a single H2O layer and basal spacing of about
~12.0 Å in Mg-bearing saponite (Suquet et al. 1975; Suquet and
Pezerat 1987; Rampe et al. in preparation), inconsistent with
observations from Cumberland. Vermiculite retains interlayer
H2O more effectively at low RH compared to saponite (Rampe et
al. in preparation), but, as previously discussed, there is no XRD
829
evidence for a vermiculite precursor in the Sheepbed member.
Based on the Ca-sulfate bearing vein content of John Klein,
Ca2+ is a plausible prediction as the dominant interlayer cation,
but Ca-bearing saponites also retain a single H2O layer (~12.0
Å spacing) at low RH values (Suquet and Pezerat 1987; Ferrage
et al. 2005; Rampe et al. in preparation). Based on the 10 Å 001
peak position and the discussions presented above that preclude
illite, monovalent interlayer cations (Na+ and/or K+) may be
prevalent in John Klein.
Our favored explanation for the lack of collapse in Cumberland is partial intercalation by metal-hydroxyl groups. On Earth,
partial intercalation or chloritization of expandable clays is most
commonly reported in acidic soils where conditions allow mobilization of hydroxy-Al that may be taken up in the interlayer
(Rich 1968). Intercalation of hydroxy-Mg groups occurs in saline
lake sediments exposed to alkaline Mg-rich waters (Jones and
Weir 1983) and via neutral to mildly alkaline hydrothermal alteration (Pevear et al. 1982). Intercalation by Fe-hydroxyl groups
is rarely observed but has been reported in reducing sediments
that were subsequently exposed and oxidized (Lynn and Whittig
1966). Partially intercalated smectites, also known as intergrade
smectite/chlorite, have a basal peak at ~13.5 Å and typically lack
a measurable 002 peak at ~7 Å. It is possible that the weak 10
Å peak in the Cumberland pattern represents a fraction that has
not been intercalated by metal-hydroxyl groups.
Origin of clay minerals
Fe/Mg-smectites are the most common types of clay minerals
observed on Mars by global spectroscopic surveys (Ehlmann et
al. 2011). The dominantly mafic composition of the martian crust
and multiple mechanisms and instances in which moderate pH
fluids can interact with abundant, reactive Fe and Mg silicate
minerals ensures this (Bristow and Milliken 2011). Thus, determining whether the clay minerals in the Sheepbed mudstone
are detrital or formed in place is challenging (Milliken and Bish
2010; Bristow and Milliken 2011).
Chemical and mineralogical data collected by MSL provide
evidence of an in situ origin for Sheepbed clay minerals. The
bulk rock chemistry of the Sheepbed member is not significantly
different from average martian basalt (McLennan et al. 2014).
Transport and hydrodynamic sorting of clay mineral-bearing
sedimentary detritus during deposition of the Sheepbed member
would likely have fractionated minerals and produced changes
in chemistry relative to other Sheepbed units, but this is not observed (McLennan et al. 2014). Instead, low water:rock alteration
in a system with little chemical transport is indicated. Consistent
with this scenario, the low forsterite content of John Klein and
Cumberland (estimated at 5.7 and 1.9 wt% of the basaltic detritus) suggests that the clay minerals in the Sheepbed mudstone
are largely derived from aqueous alteration of olivine, involving
minimal loss of chemical components (McLennan et al. 2014;
Vaniman et al. 2014). Given their sedimentary context, the high
levels of magnetite (8.7 and 9.5 wt% of crystalline phases other
than smectite for John Klein and Cumberland, respectively) are
hypothesized as a second product of the olivine alteration reaction (McLennan et al. 2014; Vaniman et al. 2014).
In the absence of active removal of chemical components,
hydration of olivine to produce clay minerals plus magnetite
830
BRISTOW ET AL.: CLAY MINERALS AT YELLOWKNIFE BAY
involves a significant increase in volume. Signs of brittle deformation expected in a lithified host were not observed in the
Sheepbed mudstone. Instead, the early diagenetic raised-ridges
observed in the Sheepbed member (Grotzinger et al. 2014) are
thought to be a response to volume changes induced by production of clay minerals prior to lithification of the Sheepbed
mudstone (Siebach et al. 2014).
An alternative hypothesis for the clay minerals at YKB is
that they are predominantly detrital materials sourced from the
northern walls and rim of Gale crater. Indeed, clay minerals have
been detected in the walls of Gale crater (Wray 2013) and in the
bedrock of the region (Ehlmann and Buz 2015). Possible modes
of detrital clay mineral formation include; (1) weathering, (2)
hydrothermal activity related to the impact that created Gale
(Schwenzer et al. 2012), and (3) regional subsurface aqueous activity prior to formation of Gale (following Ehlmann et al. 2011).
In our opinion, inputs of detrital clay minerals at YKB produced via impact-related hydrothermal alteration (Schwenzer et
al. 2012) and/or weathering in the source region were minimal.
This inference is based on comparing the quantity and types of
clay minerals in the YKB assemblage with observations and
predictions from impact-related hydrothermal systems and
weathering profiles.
An impact big enough to form a crater the size of Gale would
have induced high-temperature gradients and, given sufficient
water supply, sustained hydrothermal alteration in the crater rim
lasting up to hundreds of thousands of years (Abramov and Kring
2005). Based on terrestrial data from impacts and geochemical
modeling, a heterogeneous assemblage of clay alteration products, potentially including nontronite, celadonite, corrensite,
chlorite, mica, talc, and mixed-layer minerals composed of
these end-members could result from such processes with steep
temperature and fluid gradients (Hagerty and Newsom 2003:
Schwenzer et al. 2012). The simple clay mineral assemblage of
the Sheepbed member is not easily reconciled with such a source.
Saponite is generally produced during the early stages of
hydrolytic weathering of basalt when Fe, Mg, and Si activities
are high (Meunier 2005; but see Ziegler et al. 2003 for an exception where halloysite is the main product from inception of
arid-zone basalt weathering). However, saponite does not typically form in isolation; and the mosaic of microenvironments
within weathering profiles can generate other dioctahedral 2:1
phyllosilicates such as celadonite, Fe-beidellite and nontronite
during early basalt weathering (Meunier 2005; Garcia-Romero et
al. 2005), minerals not detected by CheMin at YKB. Moreover,
the abundance of secondary phases found at YKB, with clay
minerals making up ~20 wt% of the mudstone (Vaniman et al.
2014), suggest an extensive degree of alteration in any hypothetical weathering profile. Mature basaltic weathering profiles often
give kaolinitic clay minerals and goethite (Meunier 2005), also
not observed at YKB.
The clay mineral products of widespread Noachian subsurface
alteration of thick sections of martian crust hypothesized by
Ehlmann et al. (2011) cannot be entirely ruled out as potential
detrital source of clay minerals at YKB. However, CheMin has
not detected predicted accompanying phases such as zeolites
(Ehlmann et al. 2011), which would strengthen the case for
such a source.
In some low-temperature terrestrial sedimentary basins,
strongly saline, acidic or alkaline fluids are capable of overprinting/transforming detrital clay minerals, thus, altering clay
mineral assemblages to reflect ambient conditions prior to
lithification (Jones and Spencer 1999; Larsen 2008; Bristow et
al. 2012). The Sheepbed mudstone lacks significant quantities
of chemical phases (with the exception of later vein fills) that
would be expected in saline or acidic lacustrine systems (Bristow
and Milliken 2011; Bowen et al. 2012; McLennan et al. 2014).
Alkaline conditions (pH 9–10) could potentially explain the
partial intercalation of Cumberland smectites by metal-hydroxide
groups in addition to homogenizing clay mineral assemblage, but
the Sheepbed mudstone lacks other authigenic mineral indicators,
like zeolites, K-feldspars, and carbonates that are characteristic
of alkaline systems (Bristow and Milliken 2011).
Timing of clay mineral formation
The kinetics of olivine dissolution at low temperatures is
consistent with the alteration of olivine to a clay mineral prior
to lithification of the Sheepbed mudstone. Olivine dissolution,
the first step in the formation of secondary phases, is influenced
by various factors including grain size, pH, olivine composition,
chemistry of fluids and atmosphere, as well as temperature
(Stopar et al. 2006; Olsen and Rimstidt 2007; Hausrath and
Brantley 2010). Broad constraints are available for many of these
governing factors, including: (1) grain size <50 μm (Grotzinger et
al. 2014), (2) remnant olivine composition (estimated as Fo46 in
the Sheepbed mudstone; Vaniman et al. 2014), and (3) the persistence of clay minerals (excludes pH extremes and brackets fluid
pH in the ~5–9.5 range). The lifetime of olivine in such a system
could range from thousands to hundreds of thousands of years
at temperatures between 0 and 25 °C, when considering field
dissolution rates are typically 100× slower than laboratory rates
(Stopar et al. 2006; Olsen and Rimstidt 2007). These durations
overlap with the estimated depositional period of hundreds to
tens of thousands of years for the Sheepbed member (Grotzinger
et al. 2014). Indeed, it is plausible that the YKB formation is part
of a much thicker stack of fluvio-lacustrine strata representing
water activity over a period of millions to tens of millions of years
(Grotzinger et al. 2014), providing ample time in situ for clay
mineral production. We also note that other abundant basaltic
minerals in the Sheepbed mudstone (pyroxenes and plagioclase)
have lifetimes that are orders of magnitude greater than olivine
under surficial conditions (Hausrath et al. 2008), consistent with
their persistence in the Sheepbed member.
Clay mineral diagenesis?
XRD data suggest that YKB smectites have not undergone
detectable diagenetic changes typically associated with thermal
alteration that often promotes the transformation of saponite into
corrensite, an ordered mixed-layer trioctahedral smectite/chlorite
(Chang et al. 1986; Meunier 2005; Derkowski et al. 2013). The
ordered interstratification of smectite and chlorite yields a 001
supercell peak corresponding to a spacing of 14.2 Å (chlorite
layer) plus the spacing of the smectite layers (Moore and Reynolds 1997). Thus peaks at ~24 and ~28 Å would be present for
John Klein and Cumberland, respectively, if ordered corrensite
was present; these supercell peaks were not observed. However,
BRISTOW ET AL.: CLAY MINERALS AT YELLOWKNIFE BAY
chlorite/smectite interstratifications may be poorly ordered, and
supercell peaks are therefore not always expected products of
prograde diagenesis of saponite. In summary, minor amounts
(estimated at <5 wt% of the bulk sample because of increased
low angle scattering) of corrensite in the Sheepbed samples could
remain undetected by CheMin.
The partial intercalation of metal-hydroxyl groups suggested
by the expanded basal reflection in Cumberland is distinct from
the true mixed-layer chlorite/smectites expected from prolonged
thermally induced diagenesis. The Sheepbed member has experienced an episode of fluid induced fracturing post-lithification,
with fractures then filled by Ca-sulfates. This does not appear to
have caused the partial metal-hydroxyl intercalation of the clay
minerals in Cumberland. As discussed previously, rock permeability and porosity are likely to have been occluded during early
diagenesis (McLennan et al. 2014). Moreover, the fluids were
enriched in Ca, rather than Mg, Al, or Fe, which are necessary
to form intercalated metal-hydoxyl groups in the clay minerals. In addition, fracture density is higher in John Klein than
Cumberland, while intercalation is only apparent in Cumberland
Parts of the Sheepbed member contain early diagenetic
raised ridges (Siebach et al. 2014), which bear Mg-silicate cements tentatively assigned as Al-poor Mg-phyllosilicates based
on chemical compositional analysis by APXS and ChemCam
(McLennan et al. 2014; Léveillé et al. 2014). We propose that
the fluids from which these cements precipitated induced partial
intercalation of saponites at Cumberland. Direct precipitation of
Mg-clay minerals at low temperatures requires either high-pH
or high-Mg activities in solution (Bristow and Milliken 2011)
plus a source of aqueous SiO2; these conditions are expected to
promote partial intercalation of pre-existing smectites.
Origin of magnetite
Like saponite, magnetite is proposed to be authigenic in origin
(Vaniman et al. 2014; McLennan et al. 2014). Enrichment via
hydrodynamic sorting is excluded because overlying sandstones,
where hydrodynamic concentration of heavy minerals is expected
to be most pronounced, do not contain significant magnetite, as
indicated by their Fe content (McLennan et al. 2014). In addition,
other elemental indicators provided by APXS analysis of the
Sheepbed mudstone do not support enrichment of heavy minerals
by sedimentary processes (McLennan et al. 2014). Production
of magnetite has been tentatively tied to saponite production
during the alteration of olivine (Vaniman et al. 2014) in a process that may be analogous to aqueous alteration of olivine and
mesostasis or impact glass observed in certain chondrites and
martian meteorites (Tomeoka and Buseck 1990; Treiman et al.
1993; Keller et al. 1994).
The reaction also invites comparisons with reactions involved
in the serpentinization of ultramafic rocks on Earth, in which
olivine is hydrated, forming serpentine-group phyllosilicates,
magnetite, and hydrogen (under some conditions; see McCollom
and Bach 2009; Klein et al. 2013). Saponitic clay minerals are
not observed in typical serpentinization assemblages (Evans et
al. 2013) because their formation requires the addition of Si and
Al. In meteorites, these constituents are supplied via reaction
with mesostasis or impact glass, and it is possible that the Si- and
Al-bearing X‑ray amorphous material, ~30 wt% of the Sheepbed
831
mudstone, was a source (Morris et al. 2013).
Given the isochemical nature of the proposed clay mineral
formation reaction, mass balance between reactants and products
is expected. Thus, estimated quantities of saponite determined
from CheMin data can be used to back-calculate olivine consumed and expected quantities of magnetite co-produced. Two
scenarios are presented in Table 2. In both, it is assumed that
18 wt% saponite (the quantity observed in Cumberland) forms
with magnetite by aqueous reaction of Fo50 olivine and SiO2 and
Al2O3 from the amorphous component. In the first scenario, the
saponite has a structural formula that closely approximates the
Clay Minerals Society’s source clay saponite standard SapCa-2
[Mg3(Si3.6Al0.4)O10(OH)2]. In the second, the structural formula
is a ferrian saponite with a composition approximating material
from Griffith Park, California [Mg2Fe3+(Si3.4Al0.6)O10(OH)2; Treiman et al. 2015]. The quantities of magnetite calculated to form
with ferrian saponite are close to observed amounts, 3.4 vs. 4.4
wt% of the whole Cumberland sample, respectively. The small
difference could be accounted for by small quantities of detrital
magnetite in the Sheepbed mudstone and/or errors in mineral
quantification based on profile modeling of the XRD data. The
in situ formation of ferrian saponite requires consumption of
~15 wt% olivine, which is reasonable given orbital estimates of
basalt olivine content in the YKB/Gale crater region and olivine
content present in martian basalts and soils (Rogers and Bandfield
2009; Bish et al. 2013; Blake et al. 2013). The production of a
Fe-free saponite like SapCa-2 at YKB (reaction 1, Table 2) is
less likely given the large amounts of magnetite produced (11
wt%) by this reaction and the need for more olivine (~25 wt%).
If this “saponitization” reaction took place within the
Sheepbed mudstone it raises the possibility of H2 production
within sediments. Generation of H2 may help explain voidbearing early diagenetic concretions (“hollow nodules”) observed in the Sheepbed member (Grotzinger et al. 2014). More
importantly, H2 would have been a potential energy source for
chemolithoautotrophic organisms (Grotzinger et al. 2014; Vaniman et al. 2014).
However, studies of serpentinizing systems give rise to questions about the thermodynamic viability of producing significant
H2 from alteration of olivine to magnetite at temperatures less
than ~100 °C. Although phyllosilicate production proceeds,
thermodynamic reaction models of olivine serpentinization (at
low SiO2 activities) suggest that formation of magnetite below
100 °C is unfavorable. This limits the production of H2 because
less Fe is oxidized by water and what little Fe3+ is produced is
partitioned into Fe3+ hydroxide (McCollom et al. 2009; Klein
et al. 2013; Mayhew et al. 2013). Empirical studies of lowtemperature serpentinization reactions confirm reaction models
with Fe3+ hydroxide and ferrobrucite, the main Fe-bearing reaction products at 55 °C (Mayhew et al. 2013). These experiments
also indicate that spinel-group minerals provide an important
pathway for the oxidation of Fe by water with production of H2
(Mayhew et al. 2013).
The thermodynamic modeling approach has been adapted to
simulate hypothesized reactions at YKB. In these simulations,
fayalite was reacted with SiO2 and Al2O3, to mimic posited reactive components of the X‑ray amorphous phase, with reducing
solutions containing small amounts of sodium chloride using
832
BRISTOW ET AL.: CLAY MINERALS AT YELLOWKNIFE BAY
Geochemist Work Bench (Table 3). In this batch of simulations
anoxic conditions were prescribed because they predominate
in serpentinization systems. The influence of variable oxygen
availability is described later in this paper. End-member fayalite
was used as a reactant even though the olivine at YKB is probably Fe-forsterite (Vaniman et al. 2014), because GWB is not
optimized for dealing with solid solutions. Moreover, previous
studies indicate that magnetite forms at lower temperatures from
fayalite compared to more Mg-rich olivines (Klein et al. 2013).
The standard thermodynamic database for GWB 9.0.7 was modified to contain ferrous saponite data from Wilson et al. (2006).
The simulations suggest that the addition of reactive SiO2 and
Al2O3 does not influence the temperature at which magnetite
becomes a significant part of the alteration assemblage and H2
is produced in this type of reaction (Table 3).
Based on current knowledge, production of several weight
percent magnetite at surficial temperatures shortly after sedimentation via “saponitization” is improbable. It is possible to envisage olivine initially reacting to form saponite and Fe-hydroxides
or other poorly crystalline Fe-oxyhydroxide phases that later
matured to magnetite under more favorable higher-temperature
conditions. However, considering maximum potential burial
depth and estimated geothermal gradients, in addition to the
absence of prograde diagenetic clay minerals corrensite and
chlorite, the Sheepbed member was probably not heated above
75 °C (Vaniman et al. 2014).
Evidence of a possible low-temperature pathway comes
from studies of aqueous olivine alteration to assemblages of
saponite and magnetite in martian meteorites (e.g., Treiman et
al. 1993), although we note this style of alteration is not entirely
analogous to YKB because it only affects up to about 1% of
some SNC meteorites. For instance, in the nakhlites, there is a
succession of siderite–saponite–minor Fe oxide–saponitic gel
interpreted as post-impact hydrothermal alteration starting with
a high-temperature carbonate forming phase and cooling to
about 50 °C before forming saponite and serpentine (Changela
and Bridges 2010; Bridges and Schwenzer 2012). The Fe-rich
saponite and saponitic composition gel identified in the nakhlites
has been shown through X‑ray absorption spectroscopy to be ferric (Hicks et al. 2014), indicating oxidizing conditions. Further
exploration of the influence of different host rocks and mineral
mixtures on clay mineral formation at YKB (based on nakhlite
martian meteorites) is presented within an accompanying paper
(Bridges et al. 2015).
In the Bali CV3 chondrite, saponite and magnetite preferentially form in planar defects of olivine crystals that were subject to
oxidation and partial conversion to Fe3+-oxides during magmatic
Table 2.Mass-balance calculation for reaction between olivine and
water to form saponite and magnetite
Saponite MagnetiteFo50 olivine
observedproduced required
Reaction 1
18 wt%
11%
24.5%
Reaction 2
18 wt%
3.4%
15.1%
Note: Both reactions require balancing via addition of oxidants as reactants or
production of hydrogen gas—discussed in the text. Reactions:
(1) 2.8H2O + 0.6SiO2(aq) + 0.4Al3+(aq) + 3MgFeSiO4 → Fe3O4 + Mg3(Si3.6Al0.4)
O10(OH)2 [saponite].
(2) 2.53H 2 O + 1.4SiO 2 (aq) + 0.6Al 3+ (aq) + 2MgFeSiO 4 → 0.33Fe 3 O 4 +
Mg2Fe3+(Si3.4Al0.6)O10(OH)2 [Fe3+ saponite].
Table 3. Results of thermodynamic reaction modeling of “saponitization” under anoxic conditions at various water:rock ratios
(W:R)
Temperature
10 °C
50 °C
90 °C
130 °C
W:R = 1
Fe-saponite (wt%)
73.4
73.5
74.0
87.8
Chlorite (wt%)
5.7
5.7
5.6
2.2
Serpentine (wt%)
20.9
20.8
20.3
4.8
Magnetite (wt%)
trace
trace
0.2
5.2
H2(aq) (mmol/kg)
0.06
0.5
9.5
245
W:R = 10
Fe-saponite (wt%)
73.5
73.8
79.2
92.1
Chlorite (wt%)
5.7
5.6
4.3
1.2
Serpentine (wt%)
20.8
20.5
14.4
Magnetite (wt%)
trace
trace
2.1
6.7
H2(aq) (mmol/kg)
0.06
0.5
9.5
30.4
W:R = 100
Fe-saponite (wt%)
74.2
77
92.1
92.1
Chlorite (wt%)
5.5
4.8
1.1
1.1
Serpentine (wt%)
20.1
17
Magnetite (wt%)
0.2
1.2
6.8
6.8
H2(aq) (mmol/kg)
0.06
0.06
3
3
Notes: The production of magnetite and H2 only occurs to a significant extent
above ~75 °C. A similar temperature dependence is familiar from previous
work on serpentinizing systems (McCollom et al. 2009; Klein et al. 2013). Starting solution composition: Fe2+ = 0.1 mmol/kg, SiO2(aq) = 0.1, Cl– = 1.2 mmol/kg
(free to balance charge), Al3+ = 0.001 mmol/kg, Na+ = 1 mmol/kg, O2(aq) = 0.01
mmol/kg, pH = 7.
cooling (Keller et al. 1994). Laboratory genesis of magnetite at
room temperature and near neutral pH requires aqueous Fe2+
and small amounts of oxidants, either Fe3+ from dissolution
of ferrihydrite or air bubbled through solutions (Taylor et al.
1987). Such conditions have been likened to those found in soils
that bear enrichments of authigenic magnetite (e.g., Maher and
Taylor 1988).
A slight modification of our reaction path models that includes modest additions of O2(aq) confirms these empirical and
petrological observations (Table 4). In these models magnetite
and saponite are produced at 5 °C through moderate oxidation
of Fe2+ (Table 3). Only a trace of H2(aq) is produced and is independent of the amounts of magnetite produced. At higher [O2(aq)]
[20 mmol/kg O2(aq) using the conditions prescribed in Table 3]
hematite and gibbsite are produced at the expense of magnetite
and clay minerals and the pH at the end of the reaction drops
compared to reactions with lower [O2(aq)] (Table 4).
Terrestrial analogs
To the best of our knowledge, an exact low-temperature
terrestrial analog for the Sheepbed mudstone has not been
documented. However, assemblages of authigenic magnetite and
smectite have been reported from Eocene age iron-meromictic
Table 4.Thermodynamic reaction modeling of aqueous alteration of
fayalite, Al2O3, and SiO2, with variable amounts of O2(aq)
O2(aq) reacted (mmol/kg)
0
1
5
10
15
20
W:R = 10, T = 5 °C
Fe-saponite (wt%)
73.5 74.679.6 86.291.992.7
Chlorite (wt%)
5.7 5.44.2 2.71.2
Serpentine (wt%)
20.8
19.5
14.0
7.1
0.1
Magnetite (wt%)
trace0.52.2 4.46.73.3
Gibbsite (wt%)
0.2
Hematite (wt%)
3.8
pH
8.2 8.68.6 8.68.67.8
H2(aq) (mmol/kg)
0.05 0.050.05 0.050.05trace
Note: Starting solution composition: Fe2+ = 0.1 mmol/kg, SiO2(aq) = 0.1, Cl– = 1.2
mmol/kg (free to balance charge), Al3+ = 0.001 mmol/kg, Na+ = 1 mmol/kg, O2(aq)
= 0.01 mmol/kg, pH = 7.
BRISTOW ET AL.: CLAY MINERALS AT YELLOWKNIFE BAY
lake sediments in the Seward Peninsula, Alaska (Dickinson
1988). The basaltic flows that dammed regional drainage, resulting in lake formation, also served as a source of Fe in lake waters
and pore-fluid (Dickinson 1988). Hollow magnetite spherules
up to 0.9 mm in diameter with a proposed authigenic origin are
reported from recent sediments (<20 000 yrs old) of volcanic
crater lake Barombi Mbo in West Cameroon (Cornen et al. 1992).
Here magnetite coexists with authigenic brown granules consisting of kaolinite and Fe-hydroxides. Lake sediments are supplied
with basaltic detritus including olivine from the catchment via
streams and ash falls (Cornen et al. 1992).
Low-temperature authigenic magnetite formation within the
Sheepbed member would require moderate levels of oxidants
to react with Fe2+ derived from the dissolution of fine-grained
olivine detritus. Possible sources that could maintain such levels
include the diffusion of O2 from the atmosphere, a supply of Fe3+
from dissolution of detrital material, or photochemical production
of Fe3+ within the water column (Hurowitz et al. 2010). It is possible that the waters of YKB contained sulfate that served as an
oxidant. Oxidized and reduced forms of sulfur and perchlorates
are detected by CheMin and SAM in the Sheepbed mudstone
(Vaniman et al. 2014; Ming et al. 2014), but it is unclear as
to what proportions of these species were part of the original
depositional system as opposed to being introduced during later
diagenesis that resulted in sulfate-filled veins. Whatever the
chemical pathway, the presence of saponites indicates that the
degree of Fe2+ oxidation was moderate, and thus avoided reducing
the pH to the point where acidic conditions inhibited smectitic
clay formation and preservation was no longer favorable. Fe2+
oxidation could potentially be limited by rapid partitioning into
authigenic magnetite and clay minerals.
Implications
Geological context of the Yellowknife Bay Formation
Part of Curiosity’s ongoing mission is to determine how
the rock units the rover encounters fit into a wider geological
context of Gale crater. This is of particular importance for the
YKB formation because collective evidence indicates that the
strata represent an ancient environment potentially capable of
supporting life (Grotzinger et al. 2014). Thus, determining the
relationship between the YKB formation, the Peace Vallis fan,
and the strata at the base of Mt. Sharp have implications for its
age and help define periods in martian history when the planet
was habitable (Grotzinger et al. 2014).
On Earth, clay mineralogy is widely used as a tool for
reconstructing the burial history of sedimentary basins. Thus,
it is also worth considering whether YKB clay minerals could
provide constraints that help place the YKB strata in a regional
context. In terrestrial sedimentary basins, saponite is chloritized
to corrensite during burial diagenesis at temperatures as low
as 60–70 °C (Chang et al. 1986). However, the local chemical
environment and duration of heating have a significant influence
and conversion to corrensite may not occur until 160 °C in some
basins (Meunier 2005). The early occlusion of permeability and
porosity in the Sheepbed member may have also inhibited clay
mineral diagenesis (e.g., Derkowski et al. 2013), in combination
with dry conditions in Gale crater after clay mineral formation
833
(Tosca and Knoll 2009).
In addition to the inherent complexities of trying to use the
alteration of saponites as a paleothermometer, significant differences in geothermal gradients are expected between Earth and
Mars, given the thicker martian lithosphere and absence of plate
tectonics. A geothermal gradient of 15 °C/km is estimated for
Gale crater during the Late Noachian/Early Hesperian (Hahn et
al. 2011). The result is that a complex and prolonged history of
burial, compaction, fluid circulation, and cementation is possible
in the YKB formation with minimal heating. Indeed, comparisons
of the YKB formation with first-cycle basaltic sediments on
Earth may be particularly difficult given Earth’s active tectonic
settings and elevated geothermal gradients.
Strata at the base of Mt. Sharp may have been buried to depths
of up to 5 km (Malin and Edgett 2000), corresponding to an estimated temperature of 75 °C (Hahn et al. 2011), and would likely
have spent millions of years at this depth before being exposed
by wind erosion (Malin and Edgett 2000). Indeed, Farley et al.
(2014) calculated a surface exposure age of 78 ± 30 million years.
Such conditions would not necessarily lead to the chloritization
of saponite. Thus the absence of detectable corrensite in the
Sheepbed mudstone does not preclude a relationship between
the YKB formation and basal strata of Mt. Sharp.
It is also notable that orbitally detected clay minerals in
strata at the base of Mt. Sharp are identified as nontronite (Milliken et al. 2010), a variety of Fe-rich dioctahedral smectite
that appears to be distinct from clay minerals discovered in the
Sheepbed mudstone. Differences in clay mineralogy do not exclude the possibility that the YKB formation is part of the same
sedimentary system exposed at the base of Mt. Sharp. The clay
minerals in these two areas could simply reflect differences in
authigenic pathways influenced by local conditions or reflect
variable proportions of authigenesis vs. detrital inputs (Bristow
and Milliken 2011). Moreover, the clay minerals in Mt. Sharp
may be secondary rather than depositional or authigenic materials. Future investigations by Curiosity at Mt. Sharp will help
resolve these issues.
Mineralogic constraints on habitability
The clay minerals and authigenic magnetite discovered in
the Sheepbed mudstone provide critical paleoenvironmental
constraints. Their formation and persistence are evidence of
aqueous alteration of basaltic detritus at moderate pH (Vaniman
et al. 2014; Grotzinger et al. 2014). The additional discussion of
olivine alteration kinetics presented here indicates the persistence
of circum-neutral, moderately oxidizing aqueous conditions at
YKB for thousands to hundreds of thousands of years. This provides a first-order minimum constraint on the time available for
any life to establish itself in this ancient fluvio-lacustrine system.
The evidence presented here for low-temperature production
of saponite and magnetite by controlled oxidation of Fe2+ from
dissolved olivine also suggests potential available energy sources
for life inhabiting the muddy lake sediments of YKB. On Earth,
neutrophilic iron-oxidizing bacteria and archaea are found in a
range of habitats. Chemolithotrophic microaerophiles grow at
oxic/anoxic interfaces using soluble Fe2+ as their energy source
(Edward et al. 2004; Emerson et al. 2010), while anaerobic
organisms couple Fe2+ oxidation to nitrate reduction (Straub
834
BRISTOW ET AL.: CLAY MINERALS AT YELLOWKNIFE BAY
and Buchholz-Cleven 1998) and photoautotrophy (Widdel et al.
1993; Croal et al. 2004). Microaerophilic microbial communities are known to couple Fe2+ oxidation to Fe3+ reduction (e.g.,
rhizosphere of wetland plants; Weiss et al. 2003). Such microscale iron cycles are supported by anaerobic Fe3+ reducers such
as Geobacter metallireducens that can couple the reduction of
insoluble Fe3+ to the oxidation of a wide range of organic molecules (Lovly, et al. 1993). Perhaps most relevant to YKB are
isolates of Fe-oxidizer Pseudomonas sp. HerB from the rockice interface of a lava-tube cave in Oregon that use olivine as a
source of energy under the low oxygen conditions envisioned
within the Sheepbed mudstone (Popa et al. 2012).
H2 may have been produced through low-temperature reactions of water and olivine or in anoxic microenvironments within
the sediments at YKB (Stevens and McKinley 2000; Neubeck
et al. 2011; Mayhew et al. 2013). H2 is a viable electron donor
in a range of metabolic pathways; however, laboratory and terrestrial field observations identify H2 production rates close to
the limits of survival of archaea where water and olivine react
at low temperatures (Stevens and McKinley 2000; Neubeck et
al. 2011). Thus, H2 was, at best, a secondary potential biotic
energy source.
Acknowledgments
We are grateful to R. Kleeburg for help with modeling XRD patterns and to
D. Deocampo and J. Cuadros for thoughtful reviews. Support from the engineers
and staff of NASA Mars Science Laboratory Mission are gratefully acknowledged.
Part of this research was carried out at the Jet Propulsion Laboratory, California
Institute of Technology, under a contract with the National Aeronautics and Space
Administration.
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Manuscript received June 4, 2014
Manuscript accepted September 20, 2014
Manuscript handled by Janice Bishop